CPClimate of the PastCPClim. Past1814-9332Copernicus PublicationsGöttingen, Germany10.5194/cp-12-1445-2016Paleoclimate in continental northwestern Europe during the Eemian and early
Weichselian (125–97 ka): insights from a Belgian speleothemVansteenbergeStefsvsteenb@vub.ac.beVerheydenSophieChengHaiEdwardsR. LawrenceKeppensEddyClaeysPhilippehttps://orcid.org/0000-0002-4585-7687Earth System Science Group, Analytical-, Environmental- &
Geo-Chemistry, Vrije Universiteit Brussel, Brussels, BelgiumRoyal Belgian Institute for Natural Sciences, Brussels, BelgiumInstitute of Global Environmental Change, Xi'an Jiaotong University,
Xi'an, ChinaDepartment of Earth Sciences, University of Minnesota, Minneapolis,
USAStef Vansteenberge (svsteenb@vub.ac.be)5July2016127144514581February20168February201630May201615June2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://cp.copernicus.org/articles/12/1445/2016/cp-12-1445-2016.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/articles/12/1445/2016/cp-12-1445-2016.pdf
The last interglacial serves as an excellent time interval for
studying climate dynamics during past warm periods. Speleothems have been
successfully used for reconstructing the paleoclimate of last interglacial
continental Europe. However, all previously investigated speleothems are
restricted to southern Europe or the Alps, leaving large parts of
northwestern Europe undocumented. To better understand regional climate
changes over the past, a larger spatial coverage of European last
interglacial continental records is essential, and speleothems, because of
their ability to obtain excellent chronologies, can provide a major
contribution. Here, we present new, high-resolution data from a stalagmite
(Han-9) obtained from the Han-sur-Lesse Cave in Belgium. Han-9 formed
between 125.3 and ∼ 97 ka, with interruptions of growth
occurring at 117.3–112.9 and 106.6–103.6 ka. The speleothem was
investigated for its growth, morphology and stable isotope (δ13C and δ18O) composition. The speleothem started growing
relatively late within the last interglacial, at 125.3 ka, as other European
continental archives suggest that Eemian optimum conditions were already
present during that time. It appears that the initiation of Han-9 growth is
caused by an increase in moisture availability, linked to wetter conditions
around 125.3 ka. The δ13C and δ18O proxies
indicate a period of relatively stable conditions after 125.3 ka; however,
at 120 ka the speleothem δ18O registered the first signs of
regionally changing climate conditions, being a modification of ocean source
δ18O linked to an increase in ice volume towards the Marine
Isotope Stage (MIS) 5e–5d transition. At 117.5 ka, drastic vegetation
changes are recorded by Han-9 δ13C immediately followed by a
cessation of speleothem growth at 117.3 ka, suggesting a transition to
significantly dryer conditions. The Han-9 record covering the early
Weichselian displays larger amplitudes in both isotope proxies and changes
in stalagmite morphology, evidencing increased variability compared to the
Eemian. Stadials that appear to be analogous to those in Greenland are
recognized in Han-9, and the chronology is consistent with other European
(speleothem) records. Greenland Stadial 25 is reflected as a cold/dry period
within Han-9 stable isotope proxies, and the second interruption in
speleothem growth occurs simultaneously with Greenland Stadial 24.
Introduction
The last interglacial (LIG) period is known as the time interval before the
last glacial period during which temperatures were similar to or higher than
those of the Holocene period and present day (Otto-Bliesner et al., 2013).
In marine sediment cores, the LIG is defined as Marine Isotope Stage (MIS)
5e (Shackleton, 1969). The start and end of the MIS 5e period are
conventionally set at 130 and 116 ka, respectively, based on marine records
(e.g., Martinson et al., 1987). The expression of the LIG in
continental western Europe is defined as the Eemian, although it does not
coincide precisely with the isotopically constrained MIS 5e (Otvos,
2015). Given the ongoing debate about the LIG nomenclature,
clarification about the terms used in this manuscript is required. This
study focuses on speleothem archives; thus the terms “Eemian” and
“Weichselian” are preferred in the context of European continental
paleoclimate. The Eemian is defined as the optimum or acme LIG climate
conditions (the “sensu stricto” definition). Subsequent to the Eemian, the
Weichselian starts, with the early Weichselian in continental records
corresponding with the time equivalent of MIS 5d–5a. “Glacial inception”
is considered to be informal and only marks the Eemian–early Weichselian
transition.
The term Eemian was originally introduced by P. Harting in 1875 and was
characterized by the occurrence of warm-water mollusks in marine sediments
of the Eem River valley, near Amsterdam, the Netherlands.
Nowadays, the Eemian is mostly interpreted as an interval of warmer climate
associated with the spread of temperate mixed forests in areas with similar
vegetation to today (Kukla et al., 2002). However, the Eemian is also
known to be a diachronous unit (Kukla et al., 2002; Wohlfarth et al., 2013)
with a longer duration of up to 20 ka, from 130 to 110 ka, in
southern Europe as evidenced by pollen records (Sanchez
Goñi et al., 1999; Tzedakis et al., 2003). The LIG period exhibited
global mean temperatures (GMT) 1.5 to 2 ∘C higher than
the pre-anthropogenic average together with peak eustatic sea levels that
were between 5.5 and 9 m higher than present (Dutton and Lambeck,
2012). Therefore, the LIG gained a lot of attention from both paleoclimate
and climate-modeling communities for studying a warmer climate state and
potential future sea-level rise (Loutre et al., 2014;
Goelzer et al., 2015), even though the present-day
configuration of Earth's orbital forcing parameters is different
(Berger and Loutre, 2002). Following the Eemian, climate went into a glacial
mode known as the last glacial cycle, or the Weichselian in the western
European continental terminology, which lasted until the Holocene. A major
feature of climate variability during the last glacial is the occurrence of
millennial-scale, rapid cold–warm–cold cycles, known as Dansgaard–Oeschger
(D/O) events (Bond et al., 1993). These D/O cycles are
expressed as alternating Greenland stadial (GS) and interstadial (GIS)
phases in Greenland ice cores (Dansgaard et al., 1993;
NGRIP members, 2004), and they also have affinity with Atlantic cold
events registered in sea surface temperature proxies (McManus et al., 1994).
Some of the stadials are also associated with an increased flux of
ice-rafted debris (IRD) in the North Atlantic Ocean (McManus et
al., 1994). These events have been linked to changes in the strength and
shifts in the northwards extent of Atlantic Meridional Overturning
Circulation (AMOC) (Broecker et al., 1985). The exact cause of such
changes in the AMOC mode is still debated. Nevertheless, according to Barker
et al. (2015), it is more likely a nonlinear response of a
gradual cooling of the climate than a result of enhanced freshwater input
by iceberg calving, as previously proposed by Bond et al. (1995) and
van Kreveld et al. (2000). Likewise, continental pollen
records extracted from cores of Eifel Maar lakes (Sirocko et al.,
2005) or peat bogs in the Vosges region, France (Woillard,
1978; de Beaulieu and Reille, 1992; de Beaulieu, 2010), have
recorded changes in pollen assembly, attributed to D/O variability. So far,
up to 26 GSs have been identified in ice cores, with GS 26 recognized as
corresponding with the end of the Eemian interglacial period (NGRIP members,
2004).
(a) Location of the Han-sur-Lesse Cave site (red star) and other
records mentioned in this study: (1) Han-sur-Lesse; (2) Bourgeois-Delaunay Cave
(Couchoud et al., 2009); (3) Hölloch Cave (Moseley et al., 2015); (4)
Spannagel Cave (Hölzkamper et al., 2004); (5) Entrische Kirche Cave
(Meyer et al., 2008); (6) Corchia Cave (Drysdale et al., 2009); (7) MD04-2548
(Sanchez-Goñi et al., 2012); (8) La Grande Pile (Woillard, 1978); (9)
Eifel Maar (Sirocko et al., 2005); (10) NGRIP (NGRIP Members, 2004); (11) NEEM
(NEEM community, 2013); and (12) MD03-2664 (Irvali et al., 2012). (b)
Topographic map of the study area (source: NGI Belgium). The Han-sur-Lesse
Cave system is plotted in black. The sampling site of Han-9 is marked by the
yellow dot; the purple dot represents the location of cave monitoring by Van
Rampelbergh et al. (2014). The Réseau Sud and the Réseau Renversé are shown by the red and orange
box, respectively. For further explanation, see text. Figure adapted from Quinif (2006).
Speleothems are ideal for late Quaternary paleoclimate studies because of
their ability to obtain accurate chronologies with U/Th dating of up to
600 ka and their potential to yield high-resolution (up to seasonal scale)
paleoclimate records (Fairchild and Baker, 2012). The speleothem records
covering the Eemian and early Weichselian in Europe have provided detailed
paleoclimate reconstructions (Genty et al., 2013). Several speleothem
stable isotope proxies from Europe record optimum climatic conditions during
the Eemian (Meyer et al., 2008; Couchoud et al., 2009) and D/O climate
events during the early Weichselian (Bar-Matthews et al., 1999; Drysdale et
al., 2007; Boch et al., 2011). Yet so far, all records covering that time
period are located in southern Europe (Italy, southern France, Levant) or
the Alps. This study presents a new high-resolution speleothem dataset from
Belgium that expands the European coverage of last interglacial speleothem
archives northwards.
Earlier chronostratigraphic work on speleothem deposits and detrital
sediments within Belgian caves marked the presence of glacial–interglacial
cycles, with speleothem formation restricted to interglacial periods, when
warm and wet climates favored growth. Detrital sediments settle in colder
periods, with river deposits in cold, wet periods and reworked loams during
cold, dry periods (Quinif, 2006). From the 1980s onwards, speleothems
covering MIS 9 to 1 have been found in various Belgian caves (Bastin and
Gewelt, 1986; Gewelt and Ek, 1988). Quinif and Bastin (1994) analyzed
an Eemian flowstone from the Han-sur-Lesse Cave for its pollen composition
and demonstrated that vegetation above the cave area reflects interglacial
climate optimum conditions around 130 ± 10 ka. However, the dating of
this material contains large uncertainties related to the alpha spectrometric
methodology used. This study focuses on a recently obtained speleothem,
Han-9 from the Han-sur-Lesse Cave in southern Belgium. This stalagmite was
analyzed to better constrain (1) the chronology of Eemian optimum conditions
in Belgium and (2) the occurrence and signature of millennial-scale climate
variability (D/O) over northwestern Europe during the early Weichselian.
Descriptive image of Han-9 (a)δ13C plotted against
distance from top in mm; (b)δ18O plotted against distance from
top in mm; (c) high-resolution scan of the polished slabs; (d) interpretation
of the internal structure of the speleothem. Dashed lines: distinctive
layers; red lines: growth discontinuities; grey areas: dating samples,
where numbers refer to the samples in Table 1; yellow line: central axis
(sample axis); blue boxes: thin section locations; brown lines:
detrital material. (e) Stratigraphic log: colors indicate the presence of
dense calcite (yellow) or coarse calcite (grey); more intense color:
denser/coarser compared to lighter colored parts. The visual expression of
layering is indicated with the dashes in the log. Further description of the
log can be found in the figure and text.
Han-sur-Lesse Cave: geology and cave parameters
The Han-sur-Lesse Cave system is the largest known subterranean karst
network in Belgium, with a total length of ∼ 10 km. It is
located within the Calestienne, a SW–NE-trending superficial limestone belt of Middle
Devonian age. After deposition, these Paleozoic sediments underwent
Hercynian folding followed by Mesozoic erosion. The current hydrographic
network was established during the Neogene and Pleistocene, by erosion into
these folded belts (Quinif, 2006). The cave system was formed within
the Massif de Boine,
part of an anticline structure consisting of middle to late
Givetian reefal limestones, by a meander shortcut of the Lesse River (Fig. 1b). The thickness of the
epikarst zone above the cave is estimated to be
around 40 m.
The area of the Han-sur-Lesse Cave is located ∼ 200 km inland
at an elevation of 200 m above sea level (Fig. 1) and is marked by a maritime
climate with cool summers and mild winters. For the period 1999–2013,
average year temperature above the cave was 10.2 ∘C, and average
yearly rainfall amount 820 mm yr-1. The amount of precipitation does not
follow a seasonal distribution (Royal Meteorological Institute, RMI). The
dominant moisture source in northwestern Europe is the North Atlantic Ocean,
and this remains constant throughout the year (Gimeno, 2010). Modern δ18O of rainfall seasonally varies between -17 ‰ in
winter and -4 ‰ in summer (Van Rampelbergh et al., 2014).
The area above the cave mainly consists of C3 type vegetation with
Corylus, Fagus and Quercus trees, and as a natural reserve it has been protected from direct
human influence for over 50 years (Timperman, 1989). The Lesse River
enters the cave system at the Gouffre de Belvaux and exits at the Trou de Han approximately 24 h
later. The Han-9 stalagmite was collected within the Réseau Renversé, which is the most
distal part of the Réseau Sud (Fig. 1b). The natural connection between the
Réseau Sud and other parts of the Han-sur-Lesse Cave is fully submerged, but in 1960 an
artificial tunnel was established facilitating access (Timperman, 1989).
When water of the Lesse River is high, part of the stream is redirected
through the Réseau Sud, but it does not reach the Réseau Renversé (Bonniver, 2010). Earlier studies
have shown that cave drip waters are mostly supplied by diffuse flow through
the host rock (Bonniver, 2011; Van Rampelbergh et al., 2014). The
Han-sur-Lesse Cave is partly accessible for tourists, but because of the
difficult access the Réseau Sud is protected from any anthropogenic influence.
Temperature logging for 6 months with an interval of 2 h in the
Réseau Renversé shows an average cave temperature of 9.45 ∘C with a standard
deviation < 0.02 ∘C, which reflects the average
temperature of 9.2 ∘C above the cave for 2013. Minimum and maximum
temperatures were 9.39 and 9.51 ∘C, respectively (C. Burlet, personal communication, 2014).
This shows that the temperature in the Réseau Renversé is constant
through the year and that it reflects the average temperature above the
cave. In contrast, recent cave monitoring in more ventilated parts of the
cave indicated a temperature seasonality of 3 ∘C (Van Rampelbergh
et al., 2014). For the Réseau Renverse, there are no indications that cave morphology
changed significantly since the last interglacial. Han-9, the stalagmite
presented in this study, was deliberately sampled because it was already
broken into three parts, so no other speleothems had to be destroyed.
Although the sample was broken, it was still in situ. The candle-shaped
stalagmite has a length of 70 cm (Fig. 2c–e).
U/Th measurements of the Han-9 stalagmite (University of
Minnesota). Samples are arranged in stratigraphic order, and those in bold
indicate results acquired in 2013. U decay constants: λ238=1.55125×10-10 (Jaffey et al., 1971) and λ234=2.82206×10-6 (Cheng et al., 2013). Th decay constant: λ230=9.1705×10-6 (Cheng et al., 2013). Corrected 230Th ages
assume the initial 230Th/232Th atomic ratio of 4.4±2.2×10-6. Those are the values for a material at secular equilibrium, with
the bulk earth 232Th/238U value of 3.8. The errors are arbitrarily
assumed to be 50 %.
* δ234U = ([234U /238U]activity-1)×1000.** δ234Uinitial was calculated based on 230Th age
(T), i.e., δ234Uinitial=δ234Umeasured×eλ234×T.*** BP stands for “before present”, where the “present” is defined as
the year 1950 AD.
The Han-sur-Lesse Cave has received scientific attention in the last decades,
making it the best-understood cave system in Belgium. This includes detailed
hydrographic studies (Bonniver et al., 2010) and extended cave-monitoring
surveys (Verheyden et al., 2008; Van Rampelbergh et al., 2014), leading
to successful paleoclimate reconstructions on Holocene speleothems down to
seasonal scale (Verheyden et al., 2006, 2012, 2014;
Van Rampelbergh et al., 2015). The elaborate cave
monitoring (Fig. 1b) has provided a solid foundation for understanding the
cave system and for interpreting its paleoclimate records, even back to
∼ 120 ka.
Methods and analytical procedures
All ages were acquired by using U/Th dating at the University of Minnesota
Earth Sciences Department, Minneapolis. Nine samples were analyzed in 2013,
and an additional batch of 14 samples was dated in 2015 to further improve
the age–depth model with additional sampling locations selected based on the
preliminary age model and the stable isotope (δ13C and δ18O) data (Fig. 2d). For all U/Th analyses, 150–200 mg of speleothem
calcite was milled and analyzed with a Neptune multiple-collector plasma
source mass spectrometer (MC-ICP-MS) from Thermo Scientific. Ages were
corrected assuming an initial 230Th/232Th atomic ratio of 4.4±2.2×10-6. The age datum is 1950 CE. For additional
information about the applied method, see Edwards et al. (1987), Shen et al. (2012) and Cheng et al. (2013)
and references therein. Age–depth modeling
was carried out using the StalAge algorithm of Scholz and Hoffmann (2011).
All depths are expressed in “mm dft”, with dft being “distance from top”.
All stable isotope analysis were carried out at the Stable Isotope
Laboratory, Vrije Universiteit Brussel. A total of 1118 samples were drilled
with a Merchantek MicroMill, a computer-steered drill mounted on a
microscope. Samples were taken along the central growth axis to avoid
possible effects of evaporation during calcite deposition (Fairchild et al.,
2006). For all samples, tungsten carbide dental drill bits with a diameter
of 300 µm from Komet were used. As a function of the growth rate,
1000, 500 and 250 µm sampling resolutions were applied in order to
maintain a more or less equal resolution in the time domain. For sample
locations, see Fig. 2d. Samples were kept at 50 ∘C prior to
analysis to avoid contamination. δ13C and δ18O
isotope measurements were performed on a Perspective IRMS from Nu
Instruments, coupled to a Nucarb automated carbonate preparation system and
a minor amount (< 100) on a Kiel III device coupled to a Delta plus
XL from Thermo Scientific. Two samples of the in-house standard MAR-2(2),
which is made from Marbella limestone and has been calibrated against the
international standard NBS-19 (Friedman et al., 1982), were measured every
10 samples to correct for instrumental drift. Reported values for the
MAR-2(2) are 0.13 ‰ Vienna Pee Dee Belemnite (VPDB) for
δ18O and 3.41 ‰ VPDB for δ13C.
Analytical uncertainties in standards from individual batches were ≤ 0.05 ‰ for δ13C and ≤
0.08 ‰ for δ18O on the Nu Instruments setup.
At regular intervals, a replicate sample was measured in a different batch
to check for the reproducibility of the analytical method. Outliers were
manually detected, removed and remeasured if sufficient material was
present. In addition, six 30 µm thin sections were taken along the
growth transect (Fig. 2d).
ResultsSpeleothem morphology
Figure 2c–e show the interpretation of the internal morphology of Han-9.
In the lower part of the speleothem, below 365 mm dft, the calcite was
well-laminated alternating between thick white and slightly darker layers.
In the lower 15 mm, some fine, brown detrital laminae can be seen, although
they are confined to the very base and the lateral sides of the stalagmite.
From 430 mm dft, the calcite becomes progressively coarser and layering less
expressed with alternations between thicker parts of denser and coarser
calcite. Starting from 365 mm dft, the speleothem calcite has a very coarse
appearance and layering is almost indistinguishable. This goes on until
304 mm dft, where a first discontinuity in growth, D1, appears. This
discontinuity was identified macroscopically. After D1, 100 mm of speleothem
is characterized by alternating bands of dense and dark brown calcite with
coarser, white calcite. Between 200 and 176 mm dft, a band of dense, brown
calcite is present. Within this band, very fine laminae of red-brown
material can be observed. This band ends with a second discontinuity, D2.
Following D2, the axis of growth for the next 20 mm is tilted to the right.
The entire upper section consists of dense and dark brown calcite, with
little variation except for a coarser interval between 58 and 40 mm dft.
Besides some subtle unconformities marked by the dashed lines in Fig. 2d, no
internal layering is visible macroscopically. Thin-section locations were
chosen as representative of the typical morphologies displayed in the
stratigraphic log in Fig. 2e. Fabrics are described according
to Frisia (2015). In all thin sections, the dominant fabric of
the calcite crystals was columnar (Fig. 3a). The layering, although often
very well displayed macroscopically, was not distinguishable within the thin
sections. Variations in fabric occur between macroscopically defined
“denser” and “coarser” calcite (Fig. 2e), where the latter has smaller
columnar calcite crystals with significantly more inter-crystalline porosity
often filled with fluid inclusions (Fig. 3b) and is therefore described as
columnar open. The denser morphology has substantially larger crystals with
almost no pore space and can be defined as columnar elongated. Another type
of fabric occurs within the dark brown band of dense calcite between 200 and
176 mm dft. There, the columnar fabric is replaced with smaller, more equant
calcite crystals (Fig. 3c), which are then followed by a fine layer of brown
detrital material representing D2.
Thin section images of the Han-9. For locations, see Fig. 2d. All
pictures are taken with crossed polarized light and have the same scale, and
speleothem growth direction is upwards. (a) Thin section IIIB: (elongated)
columnar calcite fabric. (b) IIIA: open columnar fabric, with fluid
inclusions in the open voids. (c) IIA: discontinuity D2 is characterized by
the presence of brown detrital material. Also note the different fabric
underneath the hiatus.
U/Th dating
The results of U/Th dating are shown in Table 1. Dating samples are labeled
by “DAT-X”, with X representing the sample number in Fig. 2d. All ages are
displayed as “ka”. In all samples, the detrital Th content, estimated by
232Th concentration and the initial 230Th/232Th atomic ratio,
is relatively low (range: 6419–208 ppt). This leads to only minor
corrections for the 230Th age (Table 1). Errors are given as 2σ
and range between ±212 years and ±666 years. The U and Th
concentrations determined in the 2013 samples allowed reduction of the
sample size, resulting in smaller errors for the 2015 samples. From 672 to
176 mm dft, all ages are stratigraphically consistent; no age inversions
occur when taking into account the 2σ error of the U/Th ages.
Between 176 and 0 mm dft, the distribution of the ages is more chaotic, with
the occurrence of several age inversions and outliers.
Stable isotopes: δ13C and δ18O
Figure 2a and b show the results of the δ13C and δ18O analyses, plotted against sample depth in mm dft. All values are
expressed in ‰ VPDB. The δ13C varies
between -3.58 and -10.30 ‰, with an average of
-7.53 ‰. The δ18O values show smaller
variations, between -5.04 and -7.02 ‰, with an average
of -5.91 ‰. Lower-amplitude variability in both δ13C and δ18O occurs in the lower part of the stalagmite,
and larger-amplitude variations are present from ∼ 400mm dft
upwards, corresponding with distinct transitions in morphology (alternating
zones of dense, browner and coarser, whiter calcite; Fig. 2e).
Han-9 age–depth model constructed with the StalAge algorithm
(Scholz and Hoffmann, 2011). The actual age–depth model is represented by
the yellow line; the grey area marks the error (2σ). Numbers
represent the sample labels (Table 1). The brown curve displays the growth
rate. Numbers in red indicate important dates and are discussed in the text.
DiscussionAge model
The StalAge algorithm (Scholz and Hoffmann, 2011) was applied to the
individual ages in order to construct an age–depth model, displayed in Fig. 4.
It is clear that the stalagmite endured three separate growth phases and
that the discontinuities, expressed in the stalagmite morphology at 302 and
176 mm dft (Fig. 2d), correspond with two hiatuses separating these three
growth phases. In the first growth phase, all ages are in stratigraphic
order and are included within the model and the 2σ error. DAT-1 has
only limited weight in the final model, and an explanation for this is given
in the algorithm specifications (Scholz and Hoffmann, 2011). During the
modeling process, the StalAge algorithm has a step where the data are
screened for the occurrence of minor outliers and age inversions. This is
done by fitting error-weighted straight lines through subsets of three
adjacent data points. However, DAT-1 is located in the basal part of the
stalagmite, so fewer subsets of three data points can be used including
DAT-1. If DAT-1 does not fit on the error-weighted straight line created
with the adjacent data points DAT-10 and DAT-11, which is the case here, then the
error of DAT-1 will be increased and the weight of DAT-1 in the Monte Carlo
simulation for the age fitting will decrease. This results in fewer solutions
where DAT-1 is included in the Monte Carlo-simulated age models. The
occurrence of substantial changes in growth rate in the boundary areas of a
speleothem sample is recognized as a limitation of the StalAge algorithm
(Scholz and Hoffmann, 2011). Even though the three growth phases were
modeled separately with StalAge, the model does not perform well with the
start of the Hiatus 2, as DAT-19 is completely excluded. Likely, this is
again caused by the fact that DAT-19 is located in a boundary area. Here,
the stalagmite petrography shows clear evidence of a significantly decreased
growth rate after DAT-16 (110.6 ka), i.e., very dense, brownish calcite with
fine laminae (Fig. 2c and e). In complex cases, such as in this study where
multiple hiatuses occur, the simplest model is still the best. Therefore,
linear interpolation combined with good observations of changes in
petrography was applied to include DAT-19 within the age model (Fig. 4, red
line). For the third growth phase, because of the occurrence of several age
inversions, the resulting age model is unreliable. Despite the fact that
ages clearly cluster between ∼ 103 and ∼ 97 ka,
the chronology of the third growth phase is only poorly constrained, and
therefore a detailed interpretation of Han-9 in terms of paleoclimate is
limited to the first two growth phases.
The first growth phase starts at 125.34+0.78/-0.66 ka with a stable
growth rate of 0.02 mm yr-1 up to around 120.5 ka. After that, the
growth rate significantly increases, with values up to 0.15 mm yr-1. At
117.27+0.69/-1.02 ka, growth ceases and the first hiatus, H1, starts. The
hiatus lasts 4.41+1.10/-1.49 ka, and at 112.86+0.47/-0.41 ka growth
phase 2 starts. DAT-4 and DAT-5 were taken 6mm below and above the
discontinuity, and the age–depth model does not show any reason to question
the timing of H1. As for the second growth phase, growth rate remains at a
constant pace of 0.04 mm yr-1 until approximately 110.5 ka, where it
decreases to 0.006 mm yr-1. At 106.59+0.21/-0.22 ka, the second
growth phase ends. Given this age–depth model, stable isotopes were analyzed
with a temporal resolution between 100 and 0.3 years, and an average of 16
years.
Interpretation of stable isotope proxiesIsotopes deposited in isotopic equilibrium?
The best test for the presence of kinetic fractionation is to have a
reproducible record (Dorale and Liu, 2009). However, in the absence of a
second stalagmite record, Hendy tests could be performed (Hendy, 1971). The
problem here is that growth rates are rather low and the layering very fine,
so it would be hard to sample precisely in one layer. Therefore, an
additional test for correlation of δ13C and δ18O
was done by calculating the Pearson's correlation coefficient on the entire
record and on the three growth phases separately (Table 2). Yet a
correlation between δ13C and δ18O does not give
conclusive evidence for the presence of kinetic fractionation, as both
δ13C and δ18O are expected to be controlled by
climate and could therefore show positive or negative covariation (Dorale
and Liu, 2009). The Pearson's coefficients reveal that there is a clear
difference between the separate growth phases. The first growth phase, with
ρ=0.024, marks no covariation. The second growth phase, with ρ=-0.467, has a substantial degree of negative covariation, whereas the
third growth phase (ρ=0.461) has a positive covariation. The
differences between the coefficients of the separate growth phases indicate
that several processes are controlling the stable isotope variability and
that the presence or absence of covariation reflects changes in climate
conditions between the growth phases rather than the presence or absence of
equilibrium. Nevertheless, equilibrium deposition between the drip water and
recent calcite in the Han-sur-Lesse Cave has been observed by Van Rampelbergh et al. (2014).
Pearson's coefficient of correlation (ρ) calculated for the
three growth phases and for the total record.
ρNo. of measurementsGrowth phase 10.024599Growth phase 2-0.467235Growth phase 30.461284Total-0.1971118Variations in speleothem δ13C
The δ13C in Han-9 is controlled by changes in vegetation
assembly above the cave. This is deduced from the match between Han-9
δ13C and the abundance of grass pollen in the assembly of
Sirocko et al. (2005) recovered from the Eifel Maar (Fig. 1a and Fig. 5).
The agreement between Han-9 δ13C and the Eifel pollen assembly
is remarkable; increases in δ13C occur when the percentage of
grass pollen increases in the Eifel record. Similar shifts in grass/forest
vegetation have even been observed the Vosges region, France, 300 km further
south (Woillard, 1978; de Beaulieu and Reille, 1992; de Beaulieu, 2010) and
in other records from northern and central Europe (Helmens, 2014, and
references therein). Also, the δ13C of recent calcite formed
with a current forest-type vegetation above the cave is ∼-8 ‰ (Table 3). These values are similar to those
observed during the last interglacial in Han-9 (125.3–117.3 ka, Fig. 5).
Changes in δ13C of several per mill are often attributed to
changes in C3 versus C4 plants (McDermott, 2004). However, a higher
abundance of grasses does not necessarily result in an increased amount of
C4 vegetation. First of all, C4 species only make up ∼ 1 %
of the total amount of vascular plant species in northwestern Europe today
(Pyankov et al., 2010). Secondly, C4 species dominantly occur in a warmer,
tropical climate (Ehleringer et al., 1997). Finally, within the subfamily of
the Pooideae, commonly referred to as the cool-season grasses and thriving in
temperate European climate, all species use the C3 pathway (Soreng et al.,
2015). The reason why speleothem calcite tends to be enriched in 13C
when vegetation is dominated by grasses is because grasses have a smaller
biomass than trees and also the amount of soil respiration is lower, both
leading to a smaller fraction of biogenic CO2 compared to (heavier)
atmospheric CO2 within the soil (Genty et al., 2003). The similarity
between the two records does not seem to hold up after 109 ka. From here on,
δ13C becomes more depleted while the Eifel record shifts
towards an assembly dominated by grasses. It is not clear what might have
caused this depletion, but since growth rates are very low after 109 ka
(Fig. 4), it could be related to prior calcite precipitation (PCP). PCP is
known to act as a control mechanism on seasonal variations in δ13C of Han-sur-Lesse speleothems, as concluded from an elaborate
cave-monitoring study by Van Rampelbergh et al. (2014). However, the study by Van
Rampelbergh et al. (2014) was carried out on a large, tabular-shaped
stalagmite with drip water discharge rates of 300 mL min-1 and growth rates
of ∼ 1 mm yr-1, so caution is required when extrapolating these
cave-monitoring conclusions to smaller, slower-growing stalagmites in a
different part of the cave system. If PCP occurs at one site in the cave, it
does not mean that it occurs over the entire cave (Riechelmann et al.,
2011). Since no additional data on Sr and Mg are currently available for
Han-9, the presence of PCP can be neither confirmed nor rejected.
Comparison of Han-9 stalagmite with other records. The shaded blue
area marks the occurrence of the late Eemian aridity pulse (LEAP) in the
Eifel record and its equivalent in other records. (a) Sea surface temperature
(SST) reconstruction from marine core MD04-2845 (Sanchez-Goñi et al.,
2012); (b) planktonic δ18O from marine core MD03-2664 (Irvali et
al., 2012); (c) NGRIP δ18O record with indication of Greenland
stadial intervals (NGRIP members, 2004); (d) Asian monsoon reconstructions
from Dongge Cave (D3 & D4 stalagmites; Yuan et al., 2004); (e) Alpine
speleothem δ18O (TKS: Meyer et al., 2008; NALPS: Boch et al.,
2011; HöL-10: Moseley et al., 2015); (f) Eifel Maar pollen assembly
(Sirocko et al., 2005); (g) June insolation for 60∘ N (Berger and
Loutre, 1991); (h, i) Han-9 stable isotope record with a seven-point moving
average and U/Th dates; (j) paleoclimate interpretation of Han-9.
δ13C and δ18O analysis of three recent
calcite samples (RC1 to RC3) from the Réseau Renversé.
δ13Cδ18ORC1-8.19-5.73RC2-8.20-6.25RC3-7.80-6.48
To summarize, the control on Han-9 δ13C variations is the
amount of biogenic CO2 in the soil, caused by changes in the vegetation
type above the cave (forest/grasses), which is directly linked to climate.
Lower, depleted δ13C values occur during warmer and wetter
periods, when vegetation is dominated by temperate trees. Higher, enriched
δ13C values of speleothem calcite correspond with a higher
abundance of grasses above the cave during colder/dryer climate intervals.
Variations in speleothem δ18O
In midlatitude Europe, several different processes (including temperature,
amount effect and ocean source) influence speleothem δ18O
variability (McDermott, 2004). A good overview of all processes possibly
influencing the speleothem δ18O is given by Lachniet (2009).
One of the main processes acting on both precipitation δ18O and
calcite δ18O is temperature. Temperature fractionation on vapor
condensation was estimated to be around 0.6 ‰ ∘C-1 (Rozanski et al., 1992), and the temperature-dependent
fractionation between cave drip water and speleothem calcite for the
Han-sur-Lesse Cave was calculated to be -0.2 ‰ ∘C-1 (Van Rampelbergh et al., 2014). The eventual relation
between temperature and speleothem δ18O will be positive; lower
temperatures will thus lead to more negative speleothem δ18O.
The temperature control has been attributed as one of the main drivers of
δ18O fluctuations in European speleothems (Boch et al., 2011;
Wainer et al., 2013). In Han-9, this temperature control is well
expressed in growth phase 2, where more positive δ13C values,
which reflect lower temperatures through changes in vegetation, correspond
with more negative δ18O (Fig. 5). This also explains the
negative covariation as shown by the Pearson's correlation coefficient for
growth phase 2 (Table 2). It is very likely that in this record the
amount of precipitation also partly influences the δ18O signal of
the speleothem. In the tropics, variations in monsoonal strength are
regarded as the main control on speleothem δ18O via the amount
effect (Wang et al., 2001), and changes in the
amount of precipitation over time have also been considered as a driver for
variations within European speleothem δ18O records (Genty
et al., 2003; Couchoud et al., 2009, McDermott et al.,
2011). Temperature and precipitation (through the amount effect) controls
are thus expected to contribute most to the speleothem δ18O
variability, but as this record covers an interglacial–glacial transition,
other processes acting on longer timescales (i.e., fluctuations in global
ice volume) should also be considered. A significant contribution is to be
expected from the variations in δ18O of the source, i.e., the
North Atlantic Ocean, because of fluctuations in global ice volume.
Waelbroeck et al. (2002) estimated that, during MIS 5d, average global
δ18O values of ocean waters were up to 0.5 ‰
higher compared to MIS 5e.
Climate in the Belgian area between 125.3 and ∼ 97 ka125.3 ka: start of speleothem growth triggered by an increase in
moisture availability
The age model implies that Han-9 started growing at 125.3 ka (Fig. 4). Among
the recent sampling missions for Belgian LIG speleothems, this is the oldest
LIG sample found so far (S. Verheyden, personal communication, 2015). Although a flowstone
from the same Han-sur-Lesse Cave was reported to start growing at 130±10 ka (Quinif and Bastin, 1994), the accuracy of that one single
alpha spectrometric dating result can be questioned. Cave systems in Belgium
are known to be very sensitive recorders of glacial–interglacial changes,
with speleothem deposition only during interglacial intervals (Quinif,
2006). Between 131.5 and 126.5 ka, the Greenland Ice Sheet experienced
enhanced deglacial melting, and reinforced AMOC conditions were also present,
as identified from the MD04-2845 core from the Bay of Biscay (Sánchez
Goñi et al., 2012). Speleothem formation, however, did occur in the
Alps and in southern Europe before 125.3 ka (Moseley et al.,
2015; Drysdale et al., 2009). This raises the question whether or not the
start of Han-9 growth is just sample specific or if it represents a real,
albeit maybe locally confined, climate event at 125.3 ka. The SCH-5 Alpine
speleothem (Fig. 1a) was continuously deposited between 134.1±0.7
and 115.3±0.6 ka (Moseley et al., 2015). Within this record, an
increase in the δ18O proxy starting at 128.4 ka and lasting to
125.3 ka was identified as a warming phase. The warming phase recognized in
this Alpine speleothem occurs just prior to the start of the Han-9
speleothem growth, emphasizing a possible link between the warming and the
start of speleothem formation in Han-sur-Lesse. In the BDInf speleothem from
southern France (Couchoud et al., 2009) (Fig. 1a), the interval between
125.3 and 123.8 ka was identified as a period with increased rainfall
amount. A warmer/wetter period is potentially expressed in northwest
European vegetation records as well, such as the Eifel Maar record (Sirocko
et al., 2005), located only 150 km from the cave site in this study. At
125 ka, a transition of a pollen assembly consisting mainly of pioneering
Betula pollen with boreal Pinus towards an assembly significantly richer in
thermophilous, broadleaf tree pollen such as Ulmus, Quercus, Corylus and Carpinus is observed. However,
the chronology of this record was not constructed independently; the start
of the Eemian was determined by cross-correlating with the U/Th dates of the
SPA-50 Alpine speleothem record of
Holzkämper et al. (2004) and set at 127 ka. The Han-9 δ13C record at 125.3 ka has
the most negative values for the entire 125.3–117.3 ka growth period,
reaching almost -9 ‰. δ18O on the other
hand registers a decrease > 0.5 ‰ during the
first 300 years of growth (Fig. 5). The low δ13C values perhaps
demonstrate that interglacial optimum conditions were already present before
125.3 ka but that an increase in moisture availability caused by enhanced
precipitation above the cave, shown by the δ18O decrease, was
the factor needed to trigger growth of Han-9.
125–120 ka: Eemian optimum
The isotope records of Han-9 are relatively stable between 125 and
∼ 120 ka (Fig. 5). The variation of δ18O seems to
be submillennial and is largely confined to between -5.7 and
-6.3 ‰. The long-term trend, as displayed by a fitted
seven-point running average (Fig. 5), shows lower variability between 125 and
∼ 120 ka, especially when compared to younger growth periods
of Han-9 (i.e., 120–117.3 and 112.9–106.6 ka). Similar observations are
made for δ13C: submillennial variability is restricted between -7
and -8 ‰, with the exception of a positive excursion
towards -6 ‰ around 122 ka, and generally more stable
than in younger time intervals. During 125–120 ka, other paleoclimate
records display stable interglacial conditions, such as speleothems from the
Alps (Meyer et al., 2008; Moseley et al., 2015) and from Italy (Drysdale
et al., 2009), and other archives including ice cores (NEEM community, 2013)
(Fig. 5). In marine records off the Iberian Margin, the 125–119 ka period
was identified as an interval of “sustained European warmth”, following a
time of enhanced Greenland melting between 131.5 and 126.5 ka (Sánchez
Goñi et al., 2012). We therefore attribute the stability of our records
to the Eemian climate optimum persisting in the Belgian area as well. This
is also supported by the constant growth rate (Fig. 4) and the speleothem
morphology, displaying a sequence of layered calcite which does not show any
significant change over the 125–120 ka period (Fig. 2c–e).
120–117.3 ka: inception of glacial conditions
At 120 ka, an increase in δ18O of 0.5 ‰ is
observed (Fig. 5). This change in δ18O of the speleothem
corresponds with an elevated growth rate (Fig. 4) and a speleothem
morphology that becomes progressively coarser, with layers that are less
expressed (Fig. 2c and e). No major changes in the δ13C are
observed. Although both the age–depth model and the speleothem morphology
support an increase in speleothem growth rate, possibly in response to an
increase in moisture availability within the cave, there is no evidence in
the Han-9 δ18O record for an increase in precipitation. Due to
the amount effect, enhanced precipitation would cause the δ18O
signal to shift towards more negative values, which is not observed in the
record. The increase in growth rate, along with the accompanying faster CO2
degassing during speleothem formation, is known to act as a possible kinetic
control on the speleothem δ18O (Hendy, 1971; Lachniet, 2009),
yet kinetic control by fast degassing would also result in positive changes
of δ13C (Baker et al., 1997), which is not the
case here. Another possible explanation could be a temperature rise causing
the elevated δ18O signal, although such locally confined
temperature increase seems unlikely, because no other records (pollen, sea
surface temperature) support this hypothesis. Most likely, this increase in
δ18O is not related to any local climate effects but reflects a
more regional signal, which could be the increase of the source δ18O of the North Atlantic Ocean, since a rise of
0.5 ‰ is in good agreement with estimations of the source
δ18O variability over the MIS 5e–5d transition from Waelbroeck
et al. (2002). As a matter of fact, a study by Hearty et al. (2007)
combining last interglacial sea-level evolution at 15 sites around the world
shows a rapid descent towards an MIS 5d low stand at 119±2 ka. This
also favors the hypothesis that the speleothem δ18O between 120
and 117.3 ka reflects changing ocean source due to ice buildup.
A severe change in the δ13C is not observed until 117.5 ka,
where a sudden increase towards -4 ‰ occurs (Fig. 5).
This 5 ‰ change takes place within 200 years and happens
just before the first hiatus in this speleothem, suggesting that the
cessation in speleothem growth is indeed caused by a climate event. The
increase in δ13C here is believed to reflect changes in
vegetation, such as an increase in grasses resulting in lower vegetation
activity, linked to a changing (drying and/or cooling) climate. The age of
this event, 117.3 ka, stands out in other studies as well. First of all, in
the NGRIP δ18O record it falls within what is identified as
Greenland Stadial 26 (NGRIP Members, 2004). Although the signature of this
GS may not be as clear as the younger GS 25 or 24, it corresponds with the
overall decreasing trend observed in the ice δ18O and is also
recognized in the more recent NEEM ice core (NEEM community, 2013). The
global character of this climate event around 117.3 ka is evidenced by
similar changes in the North Atlantic Ocean. A high-resolution study by
Galaasen et al. (2014) has found perturbations of the δ13C of benthic foraminifera in marine sediment cores, interpreted as a
sharp decrease in North Atlantic Deep Water (NADW) formation at 116.8 ka, in
contrast with the high NADW formation observed during MIS 5e. Such
reductions in NADW led to changes in AMOC, resulting in a reduced ocean heat
transport and eventually cooling of the climate. This hypothesis is further
evidenced by lower sea surface temperatures, shown by an increase in
planktonic foram δ18O, in marine core MD03-2664 at that time
(Irvali et al., 2012) (Fig. 5). The lower resolution in the sea surface
temperature record of core MD04-2548 retrieved from the Bay of Biscay (Fig. 1a)
could explain its absence in this archive (Fig. 5). Other accurately
dated European speleothem records mark a similar climate deterioration
around the same time. Perhaps the most obvious example is the study from
Meyer et al. (2008), where a 3 to 4 ‰ drop in δ18O of four different flowstones from the Entrische Kirche Cave (Fig. 1a) is
observed between 119 and 118 ka (Fig. 5). This large drop in
speleothem δ18O is believed to be caused by a severe cooling
and was defined as the glacial inception at the cave site. Furthermore, a
subtle depletion occurs in the HÖL-10 stalagmite (Fig. 5) and was
correlated to the δ18O drop from Entrische Kirche (Moseley et
al., 2015). Closer to the Han-sur-Lesse Cave, a similar event was also
observed in the Eifel Maar record. There it was identified as the “LEAP”, or
the late Eemian aridity pulse, and defined by an increase in varve thickness,
loess and charcoal content together with a higher abundance of grass pollen
within the assembly (Sirocko et al., 2005), which would explain the increase
in δ13C of Han-9. Yet, the timing of the LEAP predates the
Han-9 event by ∼ 1 ka. Nevertheless, we suggest that both
records registered the same event and that the offset in chronology can be
caused by the tuning of the Eifel record or the uncertainty of the Han-9
age–depth model at 117.3 ka.
117.3–97 ka: stadial–interstadial changes in the early
Weichselian
After the first hiatus, growth starts again at 112.9 ka. The LEAP event in
the Eifel Maar only lasts 468 years (Sirocko et al., 2005), yet speleothem
growth does not recover immediately after the LEAP event. A similar
observation was made for the start of speleothem growth at 125.3 ka: optimum
conditions were already present before Han-9 started growing. From this
delayed growth, it appears that climate conditions need to be more favorable
(warmer/wetter) to initiate growth than to sustain growth. During the second
growth phase of the Han-9 stalagmite (112.9–106.6 ka), interesting
differences occur compared to the earlier-formed part of the speleothem.
First of all, the variation of both δ18O and δ13C
is much larger (Fig. 5). Secondly, changes in stalagmite morphology appear,
with alternations between dense, darker calcite and whiter, coarser
calcite (Fig. 2c–e). The δ13C curve of Han-9 shows a long-term
increasing trend until a maximum of -4 ‰ is reached at
110 ka. Superimposed on this trend, (sub)millennial variability ranging
between 2 and 3 ‰ is present. In contrast, within the
δ18O record a long-term trend is not as obvious, although a
minimum of -6.8 ‰ is reached between 111 and 110 ka. The
maximum in δ13C and the minimum in δ18O correspond
well with the timing of GS 25 (110.6–108.3 ka; Rasmussen et al., 2014)
observed in the NGRIP record (plotted on the GICC05modelext timescale),
implying that the stable isotopes of Han-9 reflect the temperature decrease
of the stadial, which is likely since higher δ13C is linked to
a less active vegetation cover during colder periods (i.e., more grasses) and
lower δ18O is caused by lower temperatures. The timing in the
Han-9 record is also in agreement with the GS 25 registered in the
NALPS record, which is believed to be dominantly temperature dependent (Boch et
al., 2011). The summer sea surface temperature reconstructions for marine
core MD04-2845 show a distinct decrease of ∼ 10 ∘C
at the same time (Sánchez Goñi et al., 2012). Between 112.9 and 111 ka, the variability of δ13C and δ18O in Han-9,
predating GS 25, has an inverse relationship suggesting that δ18O is mainly temperature controlled. The ice buildup effect,
displayed by the increase in δ18O between 120 and 117.3 ka, is
cancelled out by the effect of lower temperature, causing a decrease in
speleothem δ18O. It also appears that, for a general cooler
climate state, the amplitude of variability tends to increase as well,
compared to a more stable Eemian optimum (125–120 ka). The timing of the
second hiatus (106.6–103 ka) is similar to that of the occurrence of GS 24
in the NGRIP record and is also registered in the NALPS dataset from Boch et al. (2011).
However, if the hiatus has any affinity with GS 24, this raises
the question why speleothem growth stopped during GS 24 and continued during
GS 25. A plausible explanation could be that growth never fully recovered
from the GS 25, and that less favorable conditions (cooler/dryer) during the
GS 24 interval were sufficient to cease growth. This assumption is grounded
by decreased growth rate from 110 ka onwards (Fig. 4). In the Eifel Maar
core, significant changes in vegetation occur from ∼ 112 ka,
with nearly all pollen from broadleaf trees disappearing and the transition
towards a pollen assembly dominated by coniferous trees and grasses (Fig. 5). Also,
the time periods 110–108.5 ka and 106–104.5 ka are characterized
by a significant increase in loess composition and varve thickness,
indicative for dryer conditions in that area and corresponding with the GS
25 and GS 24 intervals (Sirocko et al., 2005).
Conclusions
This study highlights the potential of Belgian speleothem proxies (i.e.,
growth, morphology and stable isotopes) as recorders of regional and local
climate change over the Eemian and early Weichselian in northwestern Europe.
The start of speleothem growth occurs at 125.3 ka. At that time, however,
nearly all of the European continental records are already within the Eemian
climate optimum state. The δ18O record suggests that the
eventual trigger starting speleothem growth was most likely moisture
availability, linked to an increase in (local) precipitation at that time.
Optimum Eemian climate conditions recorded in Han-9 occurred between 125.3
and 120 ka, and the stable isotopes and speleothem morphology indicate a
relatively stable climate state. The first signs of regional changing
climate are observed in the δ18O proxy from 120 ka onwards and
are linked to a changing ocean source δ18O, caused by
increasing ice volume. The end of the Eemian (and start of the early
Weichselian) in Han-9, at 117.3 ka, is preceded by a drastic change in
vegetation activity and/or assembly that took place within 200 years,
triggered by a decrease in moisture availability linked to a drying climate.
This eventually led to cessation of speleothem growth. This event appears to
have a broad regional signature, as it is registered in other European
records as well. However, pollen records imply that temperate vegetation
seems to persist several millennia after 117.3 ka (de Beaulieu and Reille,
1992; Sirocko et al., 2005), resulting in a longer duration of the Eemian in
other records (Tzedakis et al., 2003). In addition, Han-9
also registered three stadials occurring during the early Weichselian.
During GS 26, at 117.3 ka, the end of the Eemian in Han-9 occurs. The GS 25
equivalent occurs between 110.5 and 108.5 ka as deduced from the stable
isotope proxies, and GS 24 is represented by a hiatus in Han-9 that starts at
106.6 ka. These chronologies are consistent with other European speleothem
records. For the early Weichselian, local climate appears to be more
sensitive during the early glacial conditions as the amplitude and frequency
of isotopic shifts tend to increase significantly compared to the Eemian.
Data availability
Stable isotope time series for the first two growth
phases are available online at NOAA paleoclimate database.
Acknowledgements
We thank the Domaine des Grottes de Han for allowing us to sample the
stalagmite and to carry out other fieldwork. S. Vansteenberge thanks
D. Verstraeten for the assistance with the stable isotope measurements.
P. Claeys thanks the Hercules Foundation for the upgrade of the Stable Isotope
Laboratory at VUB and the VUB Strategic Research funding.
Edited by: A. Dutton
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